Encyclopedia of Wildfires and Wildland-Urban Interface (WUI) Fires

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Slope Winds

  • Benjamin J. HatchettEmail author
  • Michael L. Kaplan
  • Nicholas J. Nauslar
  • Craig M. Smith
  • Kellen Nelson
Living reference work entry
DOI: https://doi.org/10.1007/978-3-319-51727-8_209-1



Slope winds result from the diurnal cycle of heating and cooling of the planetary boundary layer along elevated terrain. Daytime anabatic winds flow upslope, whereas at night, katabatic winds flow downslope.


Diurnal mountain winds are generated by insolation of the planetary boundary layer during the daytime and longwave emission-driven cooling during nighttime. This heating(cooling) produces horizontal temperature gradients between air above heated (cooled) topography and the air in valleys (Barry 2008). These turbulent flows represent the lower branch of a closed thermally direct circulation resulting from inclined heated or cooled planetary boundary layers that form above topography in a stratified atmosphere (Shapiro and Fedorovich 2008). Due to this thermal forcing, these winds are also commonly referred to as thermally driven winds. Diurnal mountain winds have been long studied (see Barry (2008) and Zardi and Whiteman (2012) for reviews) due to their importance in mountain environments with regards to day-to-day weather, climatological characteristics, transport of water vapor and pollutants, triggering of moist convection, agriculture, buildings, and biodiversity.

The diurnal mountain wind system is composed of three interacting wind systems occurring at multiple scales. In growing order of spatial scale, these include the slope, valley, and mountain range (Whiteman 2000). The diurnal mountain wind system is characterized by lower wind speeds (typically 1–6 m/s) and thus easily overwhelmed by synoptic-scale winds. These winds tend only to be observed during relatively quiescent synoptic weather conditions characterized by weak, anticyclonic flow aloft (Smith and Skyllingstad 2005; Stull 2017) with negligible downward momentum transfer of upper-level flow (Whiteman and Doran 1993). In most mid-latitude regions, diurnal mountain winds are best observed during the summer months when climatological high pressure dominates.

This contribution focuses on the slope wind component of the diurnal mountain wind system. Slope winds are confined to flowing up or down the slopes of a valley wall or an isolated hill. Alternative terms for these winds are anabatic (upslope) and katabatic (downslope) winds. However, these terms refer to any flows that occur up or down the topography (Zardi and Whiteman 2012). Katabatic flows also commonly describe larger scale slope flows occurring on ice sheets such as Greenland or Antarctica (van den Broeke and van Lipzig 2003; Vihma et al. 2011). This specific type of katabatic flow is influenced by Coriolis accelerations (Shapiro and Fedorovich 2008) and has little to do with fire weather; thus it is not covered in detail here. This contribution will also not cover synoptically driven downslope windstorms, or foehn winds, which have important impacts on fire weather, fire behavior, and pollutant transport.

Driving Mechanisms of Slope Winds

The interaction of solar insolation with mountain topography produces the characteristic localized circulations under quiescent synoptic conditions and weak regional pressure gradients. The thermal perturbations producing localized slope flows result from elevational differences in potential temperature and from differential heating or cooling of slopes that creates circulations in the horizontal and vertical directions (Barry 2008). The heating (cooling) of air over a slope produces a horizontal temperature and density gradient compared to adjacent ambient air. The characteristics of thermal winds include balanced antitriptic winds flowing toward low pressure under weak Coriolis accelerations and by a downslope-oriented gravity wind component. At night, katabatic winds occur when downward gravitational acceleration exceeds the upward buoyancy acceleration. The pressure gradient force is oriented away from the slope, and parcels move downslope (Stull 2017; Fig. 1a). Daytime anabatic winds are created when a pressure gradient pointing toward the slope forms when positive (upward-directed) buoyancy exceeds gravitational acceleration. The slope prevents an opposing pressure gradient force from being established, causing parcels of air to move upslope and trapping warm air along the mountainside (Stull 2017; Fig. 1b). Frictional drag, downslope momentum of advection, depth of flow, and slope inclination, curvature, and length represent additional controls on slope wind velocities (Smith and Skyllingstad 2005; Barry 2008).
Fig. 1

Schematic adapted from Barry (2008) demonstrating dynamical and thermodynamical processes influencing katabatic (a; Vdown) and anabatic (b; Vup) winds. P represents pressure, T represents temperature, and p represents atmospheric density, with subscripts denoting columns of air above the valley (1) and slope (2). H is the depth of fluid between the valley and slope. PGF stands for pressure gradient force

Surfaces of constant potential temperature (isentropes) demonstrate the flow paths of air moving adiabatically, or in the absence of heat exchange from diabatic processes. Diabatic surface sensible and radiative heating (cooling) cause isentropes to spread (tighten) above the terrain and in the free atmosphere. After shortwave solar radiative heating has ceased due to nightfall or shading by adjacent terrain, isentropes near the surface will slope downhill indicating stably stratified conditions (Fig. 2a). In the morning as slopes become heated, isentropes point upwards (Fig. 2b), eventually intersecting terrain and indicating well-mixed conditions. Boundary layer growth and the associated anabatic winds assist in mixing out the stable valley inversion (cold pool) formed in the previous night (Smith and Skyllingstad 2005).
Fig. 2

An isentropic perspective of slope winds during nighttime (a) and daytime (b). The symbol theta (θ) denotes potential temperature surfaces. Potential temperature is the temperature an unsaturated air parcel would have if brought to a reference pressure (typically 1000 hPa) through an adiabatic, reversible process. The close packing of potential temperature contours during nighttime (a) in valleys indicates stable conditions and cold pool formation (strengthened by downslope cold air drainage), whereas upward tilting potential temperature contours during daytime (b) indicates a reduction in stability that favors upward vertical motion and upslope flow

Mass flux transported by slope flows is directly proportional to thermal forcing by available sensible heat and inversely proportional to static stability (Vergeiner and Dreiseitl 1987). This implies that slope flows are strongly constrained by the availability of radiative forcing and are determined by instantaneous equilibria (Vergeiner and Dreiseitl 1987). For instance, a disruption of slope winds occurs if clouds obscure the surface from incoming solar radiation.

Further, pressure gradients will be reduced in time as air flowing downslope (upslope) warms (cools) adiabatically. In the case of neutral or unstable atmospheres, upslope winds may not develop as air parcels will ascend vertically rather than moving along the slope (Orville 1964).

Theoretical and numerical models as well as observations indicate slope flows are nonsteady (Hunt et al. 2003; Fedorovich and Shapiro 2009; Zardi and Serafin 2015). Periodicity in temperature and wind speeds vary as a function of slope angle (Fleagle 1950), the temperature difference between the katabatic layer and the adjacent valley (McNider 1982), and horizontal advection and atmospheric stability (Zardi and Serafin 2015). Characteristic periods of slope flows are approximately 1.5 h (Doran and Horst 1981), with the period becoming shorter as stability increases (McNider 1982).

Differences Between Anabatic and Katabatic Winds

Anabatic winds tend to have higher velocities compared to nocturnal katabatic winds. This is because daytime radiative flux exchanges are greater, promoting more significant buoyancy perturbations. Through the interaction with valley wind systems and the transfer of valley air to slopes, anabatic winds will have deeper boundary layer penetration compared to katabatic winds, which are very shallow. Katabatic winds grow in depth as slope length increases with measured depths approximately 150 m above ground level (Whiteman and Zhong 2008). Typically, the height of the katabatic jet (velocity maxima) is a few meters above ground level, which places severe constraints on the ability of numerical weather prediction models to accurately simulate this phenomenon (Smith and Porte-Agel 2014).

Contribution of Katabatic Winds to Cold Pool Formation

Katabatic winds are also known as drainage flows. While small-scale cold air drainage flows and katabatic winds both result from density perturbations produced by radiative cooling, katabatic winds require a minimum slope and will also establish compensating currents (Lawrence 1954). Nonetheless, both katabatic winds and small-scale cold air drainage contribute to cold air pooling in mountain valleys and strengthen the formation of temperature inversions produced by localized radiative cooling (Fig. 2a; Smith and Skyllingstad 2005). Cold pool formation is important with regards to wildland and prescribed fires because of the impacts of smoke transport and smoke persistence on populated areas in valley bottoms (Goodrick et al. 2013) and reductions in fire activity (Sharples 2009). Katabatic winds traveling downslope spread out horizontally at the elevation where the buoyancy of the flow equals that of the stratified air in the valley (Stull 2017). Above the characteristic top of the inversion, a region of climatologically warmer minimum temperatures known as the thermal belt occurs (Fig. 2a). The warmer minimum temperatures result in lower relative humidity and greater evaporative demands that cause fuels to cure more quickly following precipitation. Due to the poor nocturnal moisture recovery and drier fuels, more active fire conditions are common along the thermal belt compared to adjacent terrain (Sharples 2009).

Environmental Influences on Slope Winds

The characteristics of diurnal mountain winds, such as their velocity, depth, duration, and timing, varies as a function of local terrain characteristics (i.e., inclination (slope) and orientation (aspect); Whiteman et al. 1989; Barry 2008; Fernando et al. 2015). Land surface characteristics such as soil moisture, vegetation, terrain orientation (aspect), as well as the underlying geology and geomorphology, influence both surface albedo and surface roughness. These surficial characteristics ultimately impact the surface energy budget and determine how air will flow up and down mountain slopes (Barry 2008). These factors may also have a seasonal dependence that affects the amplitude of horizontal pressure gradients that produce winds (Cogliati and Mazzeo 2006).

Slope flows quickly respond to variations in sensible heat flux that result from variations in the local energy budget. Clear sky conditions, which frequently coincide with anticyclonic upper level flow regimes, facilitate optimal longwave emission and surface cooling during the night and incoming solar radiation during the day. Other factors influencing net radiation on a slope include shading, elevation, soil moisture, the partitioning of latent and sensible heat fluxes, snow cover, and vegetation (Barry 2008). Correctly forecasting the timing and magnitude of slope flows in complex terrain will require improvements in both observations, data assimilation, and physical representation in numerical models of these processes (Fernando et al. 2015).

Interactions Between Slope, Valley, and Mountain-Plains Winds

Differentialheating of mountain valley sidewalls produces contrasts in instantaneous latent and sensible heat fluxes (Whiteman et al. 1989). This differential heating acts in concert with local slope flows to produce cross-valley, along-valley, and mountain-plain wind circulations (Whiteman 2000). The anabatic (katabatic) wind is the ascending (descending) component of the cross-valley circulation, which blows toward the valley sidewall experiencing peak heating. Along-valley circulations develop at night as cold descending air flows down the valley. Warm, ascending air flows up the valley during the day. The weak low pressure that results in the valley from the removal of air as it moves upslope facilitates the formation of an upstream flowing along-valley circulation. Weak, return flow aloft will develop, oriented down-valley during the daytime and up-valley during the nighttime; these are called anti-valley winds (Stull 2017). At the mountain range scale, the mountain-plain wind system is a larger version of diurnal slope winds (Koch et al. 2001). It originates due to thermal contrasts in the air over an entire mountain range and adjacent plain with winds blowing up (down) the mountain flank during the day (night). Elevated convergence induced by thermally-forced winds at all spatial scales can facilitate moist convection (Koch et al. 2001) leading to potential fire ignitions from lightning.

An example of the diurnal variability of wind and likely interaction of diurnal wind circulations is shown in Fig. 3 for the Pleasant Grove, Utah Remote Automated Weather Station (RAWS). The data shown is for a 5-day period of calm synoptic conditions in August 2018. Pleasant Grove is located on a mid-slope along the western front of the Wasatch Mountains at the mouth of the American Fork canyon and adjacent to the Provo-Orem metropolitan area. During calm ambient weather conditions with high solar radiation in August, a cyclical pattern of easterly downslope and down-valley wind occurs during the nocturnal hours (Fig. 3a). As solar radiation increases, winds shift to westerly, indicative of the establishment of upslope and up-valley conditions. Temperatures during this period follow a strong diurnal cycle (Fig. 3b), with relative humidity typically increasing during nighttime hours. The patterns of winds observed at a weather station will tend to capture interacting slope and valley wind circulations, rather than singular circulations.
Fig. 3

Wind speed, direction, and solar radiation (a) and temperature and relative humidity (b) at the Pleasant Grove, Utah Remote Automated Weather Station (RAWS)

Applications to Fire Management

Correctly forecasting slope winds has practical applications for management of both wild and prescribed fire. Slope winds and cold air drainage can influence the distribution of air temperature in complex terrain, particularly through the formation of cold air pools. Cold pool formation, as well as the advection of temperature and moisture, influence nocturnal relative humidity recovery and fuel moisture with potential implications for fire behavior (Sharples 2009). Fires burn uphill burn faster and more intensely than those on flat slopes due to changes in flame dynamics and induced flow as well as flame proximity or contact with upslope vegetation that preheats fuels (Morandini et al. 2018). The magnitudes of slope winds (typically <5 m/s or 18 km/h) are sufficient to aid fire spread but unlikely to be the sole driver of extreme fire behavior. In the presence of stronger synoptic winds, slope flows could contribute to potential extreme fire behavior under conditions of highly receptive fuel beds. Suppression efforts must recognize the potential for rapid reversal of wind direction and thus fire growth during the morning and evening transition times (Sharples 2009; Fig. 3a). For example, a transition in wind direction could pose hazards to fire suppression personnel and values at risk if the direction of fire growth or ember-fall shifts toward unburned slopes or landscape features such as narrow canyons where flow channeling could occur (Whiteman and Doran 1993).

An important application of slope winds pertains to smoke transport and air pollution management, especially during prescribed fire operations. Under conditions of low fuel availability or connectivity, and/or high fuel moisture content, a fire may not attain sufficient energy release to create a convective column of air. In this scenario, surface winds, and not winds aloft, will bear the primary responsibility for smoke dispersal. During the smoldering phase of burning when a convective plume fails to develop, surface winds are the primary determinant of smoke transport (Whiteman 2000). Anabatic winds can blow smoke upslope and facilitate entrainment of smoke into the larger scale flow aloft, enabling long-range transport in the absence of a convective plume. Katabatic winds transport smoke downslope and promote smoke accumulation in valleys in conjunction with the cold air pool formation. This concentration of pollutants in valleys can lead to air quality concerns with regards to health and visibility (Goodrick et al. 2013). If the smoke impacts populated areas, it also may create public opposition to future prescribed burn operations (McCaffrey 2006). Careful assessment of meteorological conditions favoring or inhibiting slope winds during prescribed burns can help minimize smoke impacts on populated regions (Larkin et al. 2009) and estimate fire behavior (Sharples 2009) when wildland fires are ignited under quiescent synoptic flow regimes from dry lightning or human ignition sources. Smoke transport model assumptions need to be considered when slope winds are of concern. For instance, slope winds can violate assumptions made in simplistic Gaussian plume models (which assume steady-state, homogenous transport of pollutants), reducing their reliability (Goodrick et al. 2013). More advanced pollutant dispersion models are required if slope winds are of concern for smoke management, with improved results produced by coupled weather, dispersion, emissions, fuels, and consumption models such as BlueSky (Larkin et al. 2009).

A primary challenge in simulating slope winds results from the nearly infinite suite of potential configurations and spatial scales of terrain, geology, and vegetation in mountain environments. This limits the applicability of specific field measurements and numerical experiments to a broader understanding of the diurnal mountain wind system (Zardi and Whiteman 2012). Many numerical and analytical models have been developed and have yet to match observations, making selection of suitable models for varying circumstances difficult (Whiteman and Zhong 2008). Katabatic flows do not follow Monin-Obhukov scaling used in numerical weather prediction models (Grisogono et al. 2007). The vertical scales of motion in katabatic flows are particularly sensitive to model configuration and thus very difficult to accurately simulate (Smith and Porte-Agel 2014).


Slope winds are thermally driven winds occurring in areas of mountainous terrain and characterized by a morning and evening reversal of downslope (katabatic) and upslope (anabatic) flow. Slope winds most frequently occur during quiescent large-scale weather with clear sky conditions and are a component of the regional mountain valley-wind circulation. Although slope winds are generally shallow (near-surface) and of weak magnitude (<5 m/s or 18 km/h), they can influence fire behavior and play roles in the dispersal and transport of smoke. Fire managers are recommended to consider the timing and direction of slope winds when planning fire operations to mitigate the risks and impacts associated with wildland and prescribed fire.



  1. Barry RG (2008) Mountain weather and climate. Cambridge University Press, Cambridge, UKCrossRefGoogle Scholar
  2. Cogliati MG, Mazzeo NA (2006) Air flow analysis in the upper Río Negro Valley (Argentina). Atmos Res 80(4):263–279.  https://doi.org/10.1016/j.atmosres.2005.09.005CrossRefGoogle Scholar
  3. Doran JC, Horst TW (1981) Velocity and temperature oscillations in drainage winds. J Appl Meteorol 20:361–364CrossRefGoogle Scholar
  4. Fedorovich E, Shapiro A (2009) Structure of numerically simulated katabatic and anabatic flows along steep slopes. Acta Geophys 57(4):981–1010.  https://doi.org/10.2478/s11600-009-0027-4CrossRefGoogle Scholar
  5. Fernando HJ, Coauthors (2015) The MATERHORN: unraveling the intricacies of mountain weather. Bull Amer Meteor Soc 96:1945–1967.  https://doi.org/10.1175/BAMS-D-13-00131.1CrossRefGoogle Scholar
  6. Fleagle RG (1950) A theory of air drainage. J Met 7:277–232MathSciNetGoogle Scholar
  7. Goodrick SL, Achtemeier GL, Larkin NK, Liu Y, Strand TM (2013) Modelling smoke transport from wildland fires: a review. Int J Wildland Fire 22(1):83–94.  https://doi.org/10.1071/WF11116CrossRefGoogle Scholar
  8. Grisogono B, Kraljevi’c L, Jeriˇcevi’c A (2007) The low-level katabatic jet height versus Monin–Obukhov height. Quar J Roy Meteorol Soc 133:2133–2136CrossRefGoogle Scholar
  9. Hunt JCR, Fernando HJS, Princevac M (2003) Unsteady thermally driven flows on gentle slopes. J Atmos Sci 60:2169–2182CrossRefGoogle Scholar
  10. Koch SE, Zhang F, Kaplan ML, Lin YL, Weglarz R, Trexler CM (2001) Numerical simulations of a gravity wave event over CCOPE. Part III: the role of a mountain-plains solenoid in the generation of the second wave episode. Mon Weather Rev 129:909–933.  https://doi.org/10.1175/1520-0493(2001)129<0909:NSOAGW>2.0.CO;2CrossRefGoogle Scholar
  11. Larkin NK, O’Neill S, Solomon R, Raffuse S, Strand T, Sullivan DC, Krull C, Rorig M, Peterson J, Ferguson S (2009) The BlueSky smoke modeling framework. Int J Wildland Fire 18:906–920.  https://doi.org/10.1071/WF07086CrossRefGoogle Scholar
  12. Lawrence EN (1954) Nocturnal winds. Prof Notes Met Office (London) 7(11):1–13Google Scholar
  13. McCaffrey S (2006) Prescribed fire: what influences public approval? In: Dickinson MB (ed) Fire in eastern oak forests: delivering science to land managers. Newtown Square: USDA Forest Service, Northern Research Station. Gen Tech Rep NRS-P-1. Available at: https://www.nrs.fs.fed.us/pubs/gtr/gtr_nrs-p1/mccaffrey_p1_192.pdf
  14. McNider RT (1982) A note on velocity fluctuations in drainage flows. J Atmos Sci 39:1658–1660CrossRefGoogle Scholar
  15. Morandini F, Silvani X, Dupoy JL, Susset A (2018) Fire spread across a sloping fuel bed: flame dynamics and heat transfers. Combust Flame 190:158–170.  https://doi.org/10.1016/j.combustflame.2017.11.025CrossRefGoogle Scholar
  16. Orville HD (1964) On mountain upslope winds. J Atmos Sci 6:622–633CrossRefGoogle Scholar
  17. Shapiro A, Fedorovich E (2008) Coriolis effects in homogeneous and inhomogeneous katabatic flows. Quar J Roy Meteor Soc 134:353–370.  https://doi.org/10.1002/qj.217CrossRefGoogle Scholar
  18. Sharples JJ (2009) An overview of mountain meteorological effects relevant to fire behavior and bushfire risk. Int J Wildland Fire 18:737–754.  https://doi.org/10.1071/WF08041CrossRefGoogle Scholar
  19. Smith CM, Porte-Agel F (2014) An intercomparison of subgrid models for large eddy simulation of katabatic flows. Quart J Roy Meteor Soc 140(681):1294–1303.  https://doi.org/10.1002/qj.2212CrossRefGoogle Scholar
  20. Smith CM, Skyllingstad ED (2005) Numerical simulation of katabatic flow with changing slope angle. Mon Weather Rev 133:3065–3080.  https://doi.org/10.1175/MWR2982.1CrossRefGoogle Scholar
  21. Stull R (2017) Practical meteorology: an algebra-based survey of atmospheric science -version 1.02b. University of British Columbia, Vancouver, 940 p. Available at: https://www.eoas.ubc.ca/books/Practical_Meteorology/prmet102/Ch17-Regional-v102.pdf
  22. van den Broeke MR, van Lipzig NP (2003) Factors controlling the near-surface wind field in Antarctica. Mon Weather Rev 131:733–743.  https://doi.org/10.1175/1520-0493(2003)131<0733:FCTNSW>2.0.CO;2CrossRefGoogle Scholar
  23. Vergeiner I, Dreiseitl E (1987) Valley winds and slope winds – observations and elementary thoughts. Meteorol Atmos Phys 36:264–286CrossRefGoogle Scholar
  24. Vihma T, Tuovinen E, Savijärvi H (2011) Interaction of katabatic winds and near-surface temperatures in the Antarctic. Climate Dynam 116.  https://doi.org/10.1029/2010JD014917
  25. Whiteman CD (2000) Mountain meteorology. Oxford University Press, Oxford. 355 ppGoogle Scholar
  26. Whiteman CD, Doran JC (1993) The relationship between overlying synoptic-scale flows and winds within a valley. J Appl Meteorol 32:1669–1682.  https://doi.org/10.1175/1520-0450(1993)032<1669:TRBOSS>2.0.CO;2CrossRefGoogle Scholar
  27. Whiteman CD, Zhong S (2008) Downslope flows on a low-angle slope and their interactions with valley inversions. Part I: observations. J Appl Meteor Climatol 47:2023–2038.  https://doi.org/10.1175/2007JAMC1669.1CrossRefGoogle Scholar
  28. Whiteman CD, Allwine KJ, Fritschen LJ, Orgill MM, Simpson JR (1989) Deep valley radiation and surface energy budget microclimates. Part II: Energy Budget J Appl Meteor 28:427–437.  https://doi.org/10.1175/1520-0450%281989%29028%3C0427%3ADVRASE%3E2.0.CO%3B2CrossRefGoogle Scholar
  29. Zardi D, Serafin S (2015) Notes and Correspondence: An analytic solution for time-periodic thermally driven slope flows. Quart J Roy Met Soc 141:1968–1974.  https://doi.org/10.1002/qj.2485CrossRefGoogle Scholar
  30. Zardi D, Whiteman CD (2012) Diurnal mountain wind systems. In: Chow FK, DeWekker SFJ, Snyder B (eds) Mountain weather research and forecasting. Recent progress and current challenges. Springer atmospheric sciences. Springer, Berlin, pp 37–122Google Scholar

Copyright information

© Springer Nature Switzerland AG 2019

Authors and Affiliations

  • Benjamin J. Hatchett
    • 1
    • 2
    Email author
  • Michael L. Kaplan
    • 2
  • Nicholas J. Nauslar
    • 2
  • Craig M. Smith
    • 2
  • Kellen Nelson
    • 2
  1. 1.Western Regional Climate CenterRenoUSA
  2. 2.Division of Atmospheric SciencesDesert Research InstituteRenoUSA

Section editors and affiliations

  • Kuibin Zhou
    • 1
  1. 1.Nanjing Tech UniversityNanjingChina